Journal of the Geological Society, London, Vol. 154, 1997, pp. 73–78, 2 figs, 2 tables. Printed in Great Britain
The recognition of reactivation during continental deformation
R. E. HOLDSWORTH 1 , C. A. BUTLER 1,3 & A. M. ROBERTS 2
1
Department of Geological Sciences, University of Durham, Durham DH1 3LE, UK
2
Badley Earth Sciences Ltd, North Beck House, Hundleby, Spilsby, Lincolnshire PE23 5NB, UK
3
Present address: Elf Caledonia Ltd, Bridge of Don, Aberdeen AB23 8GB, UK
Abstract: Reactivation involves the accommodation of geologically separable displacement events
(intervals >1 Ma) along pre-existing structures. The definition of a significant period of quiescence is
central to this phenomenological definition and the duration of the interval chosen represents the
resolution limit of reactivation criteria found in most ancient settings. In neotectonic environments,
reactivation can be further defined as the accommodation of displacements along structures that formed
prior to the onset of the current tectonic regime. This mechanistic definition cannot always be applied to
ancient settings due to the uncertainties in constraining relative plate motion vectors. Four sets of criteria
may be used to recognize reactivation in the geological record: stratigraphic, structural, geochronological
and neotectonic. Some structural criteria may not be reliable if used in isolation to identify reactivated
structures. Much of the previously published evidence cited to invoke structural inheritance is equivocal
as it uses similarities in trend, dip or three-dimensional shape of structures. Numerous fault and shear zone
processes can cause significant weakening both synchronously with deformation and in the long-term and
may be invoked to explain reactivation. The collage of fault-bounded blocks forming most continents
therefore carries a long-term architecture of inheritance which can explain much of the observed
complexity of continental deformation zones.
Keywords: reactivation, faults, shear zones, deformation, rheology.
The deformation of much of the Earth’s lithosphere is characteristically heterogeneous. Strain, on all scales, is generally
focused into faults and shear zones that bound units of less
deformed material. Very important differences exist, however,
in the distribution and complexity of deformation within the
oceanic and continental regions (Molnar 1988 and references
therein). Oceanic lithosphere appears to behave in an approximately rigid manner, as required by the plate tectonic model,
in which most deformation is restricted to narrow (<100 km
wide) belts of deformation around the plate margins, with little
or no strain in the internal parts of the plate. In contrast,
continental lithosphere is characterised by broad and diffuse
zones in which deformation commonly occurs across belts
many hundreds or thousands of kilometres wide. These zones
typically comprise regions in which fault- and shear-zonebounded blocks partition strains into a series of complex
displacements, internal strains and rotations in response to
far-field plate tectonic stresses and large-scale body forces
(Dewey et al. 1986). This behaviour reflects the weakness of
continental lithosphere which is attributed to the relative
buoyancy and weak quartzofeldspathic rheology of the continental crust (e.g. Thatcher 1995). In addition, important lateral
strength variations occur due to the presence of pre-existing
structures in the continental crust such as old faults and shear
zones. These long-lived zones of weakness tend to reactivate
repeatedly, accommodating successive crustal strains, often in
preference to the formation of new zones of displacement. As
continental crust is not normally subducted, successive phases
of deformation and accretion therefore impart a long lived and
complex architecture of inheritance that is not preserved in
younger oceanic lithosphere (Sutton & Watson 1986).
Reactivation is important because pre-existing structures in
the continental lithosphere are known to strongly influence the
location and architecture of a broad range of geological
features such as orogenic belts, fault-controlled sedimentary
basins and the rifting of continents (Dewey et al. 1986; Daly
et al. 1989). Furthermore, many long-lived structures act as
channelways for the upward migration of hydrous fluids and
magmas, so that they may also largely determine the siting
and mode of emplacement of metalliferous ore deposits and
igneous intrusions (e.g. O’Driscoll 1986; Hutton 1988).
This article aims to define some basic terminology for
reactivated structures. Reliable criteria that may be used to
identify reactivation are set out, citing important examples,
and less useful, equivocal criteria are also discussed. Finally,
causative fault zone weakening mechanisms are briefly
reviewed.
Reactivation
Reactivation is here defined as the accommodation of geologically separable displacement events (intervals >1 Ma) along
pre-existing structures. These discontinuities may include
faults, shear zones, major compositional/rheological boundaries and magma ascent pathways. Reactivated structures may
display different senses of relative displacement for successive
events (geometric reactivation; Fig. 1a) or similar senses of
relative displacement for successive events (kinematic reactivation; Fig. 1b). The notion of a significant period of inactivity is
central to the meaning of a reactivated structure. In ancient
settings, most geological criteria cannot separate events where
the cessations of movement are less than 1 Ma. The shorter
term, successive displacements that occur along all active faults
as part of the seismic cycle (recurrent movements) are not
usually preserved as they typically display time intervals of
between 103 and 105 years (e.g. Wallace 1984).
In neotectonic settings, the phenomenological definition of
reactivation given here may be insufficient, since the resolution
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R. E. HOLDSWORTH ET AL.
criteria need to be recognized in order to identify reactivation
with certainty, preferably with some direct geochronological or
indirect stratigraphic evidence constraining the absolute age of
fault movements. Used in isolation, changes in the sense or
direction of movement along faults and shear zones are not
always reliable reactivation criteria. Multiple slip vectors can
arise due to the reorientation of local incremental strain and
stress fields due to slip on nearby faults (Cashman & Ellis
1994) or due to kinematic partitioning of crustal strains (Tikoff
& Teyssier 1994).
Geometric similarity: an equivocal reactivation criterion?
Fig. 1. (a) Geometric and (b) kinematic reactivation. The numbers
indicate the sequence of movement.
of displacement events has much greater precision (e.g. see
Stewart & Hancock 1994). In such cases, reactivation can be
further defined as the accommodation of displacements along
structures that formed prior to the onset of the current tectonic
regime (cf. Muir Wood & Mallard 1992). In ancient settings,
this mechanistic definition is less reliable since relative plate
motion vectors are more difficult to constrain and very similar
movements can be repeated during widely separated time
periods (e.g. as occurred in N America and NW Europe during
the opening of the modern Atlantic Ocean).
Reactivation criteria
From the existing literature, we have identified four groups of
generally reliable criteria that may be used to recognize
reactivation: stratigraphic, structural, geochronological, and
neotectonic (Fig. 2a–d). Table 1 summarizes these groups,
citing literature on important examples or reviews of each
phenomenon.
The recognition of reactivation requires evidence for repetition of displacement and associated deformation using absolute or relative time markers. Wherever possible, several
The criteria outlined in Fig. 2 and Table 1 constitute reliable
geological evidence for reactivation. In the literature (reviewed
by Prucha 1992 and Hoppin 1995), however, a number of
other indicative features have been suggested, including:
• parallelism of structures in younger rocks with those
in adjacent basement regions (e.g. Hoppin & Palmquist
1965);
• orientations of younger structures that do not match far
field stress patterns related to causative plate motions (e.g.
Donath 1962);
• parallelism of trends in the cover with geological or geophysical ‘lineaments’ that are of possible deep crustal origin
(e.g. Sutton & Watson 1986) or related to enigmatic
fracture patterns defined on a global scale (e.g. Sonder
1956; O’Driscoll 1980).
Most of these suggestions are based on apparent similarities in trend, dip or three-dimensional shape of structures, a
group of criteria that we term here geometric similarity. This
criterion is often applied to structural interpretations of seismic reflection data from offshore sedimentary basins where the
actual fault zones and underlying basement are not exposed.
Numerous publications (e.g. Brewer & Smythe 1984; BIRPS &
ECORS 1986; Gage & Doré 1986; Bartholomew et al. 1993)
concerning the structural evolution of the offshore basins
developed around the British Isles either implicitly or explicitly
assume that reactivation of underlying basement structures has
controlled faulting patterns and basin architecture. For
example, Lee & Hwang (1993) investigated the structure of the
East Shetland Basin in the northern North Sea. They recognised four structural trends, as delineated by the strike of Late
Jurassic extensional and transfer faults. The four trends (N–S,
NE–SW, NW–SE and E–W) are attributed by these authors
to basement control respectively by Archaean structures or
Silurian extensional faults; Caledonian thrust faults; Tornquist
Zone structures; and Carboniferous faults. These interpretations are based solely on the similarity in trend of each set of
structures with those in adjacent onshore basement areas. It is
likely, however, that some or all of the faults are new structures formed during lithospheric extension (Roberts et al.
1995) and, as the large amount of seismic data available from
this region provides no reliable evidence either for or against
reactivation, the model is therefore purely conjectural. We
accept that basement control of faulting in offshore basins may
be of great importance, but suggest that widespread evidence
to support this hypothesis has yet to be found in many areas
(cf. some onshore regions, Prucha 1992). Geometric similarity
cannot be used as unequivocal evidence for reactivation in the
absence of additional independent evidence. That evidence can
be acquired from detailed seismic data by constructing fault
displacement contour diagrams and looking for discontinuities
across regional unconformities (e.g. Clausen et al. 1994, fig. 6).
RECOGNITION OF REACTIVATION
Fig. 2. Reactivation criteria; refer to Table 1 for explanation and references. The isotopic ages shown in (c) are for illustrative purposes only.
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R. E. HOLDSWORTH ET AL.
Table 1. Structural inheritance criteria with examples from the literature (see Fig. 2a–d)
Inheritance criteria
(a) Stratigraphic criteria
(i)
Repeated changes in sediment package thicknesses across faults
(Fig. 2ai)
(ii)
Repeated footwall uplift unconformities (Fig. 2aii)
(iii)
Basin inversion geometries (Fig. 2aiii)
(iv)
(v)
Repeated syn-sedimentary deformation episodes (Fig. 2aiv)
Reactivation of basement faults across unconformities (Fig. 2av)
(vi)
Indirect stratigraphic evidence (Fig. 2avi)
(b) Structural criteria
(i)
Changes in kinematic history indicated by overprinting structures
(Fig. 2bi)
(ii)
Changes in distribution and nature of deformation products
within fault and/or shear zones (Fig. 2bii)
(c) Geochronological criteria
(i)
Direct dating of deformation products (Fig. 2ci)
(ii)
Indirect evidence using dated cross-cutting units (Fig. 2cii)
(d) Neotectonic criteria
(i)
Modern/historical seismicity along ancient faults (Fig. 2di)
(ii)
Offsets of geomorphological/anthropogenic features across preexisting fault trace at surface (Fig. 2dii)
Alternatively, it may be possible to trace basin-bounding faults
onshore to look for reliable reactivation criteria (e.g. see Butler
et al. 1995).
Controls of reactivation
A large number of theoretical, experimental and microstructural studies have shown that there are numerous fault and
shear zone processes which may lead to weakening both
Table 2. Processes in faults and shear zones that lead to transient and
long-term weakening
(1) Processes leading to syn-tectonic and long-term weakening
Generation of pre-existing (cohesionless) fractures
Grain refinement processes (especially grain size reduction)
General reaction softening/weakening
Geometric and fabric softening/weakening
Thermal perturbations
(2) Processes leading to transient weakening (syn-tectonic only)
Shear heating
Increases in pore fluid pressure
Transient fine-grained reaction products
Transformational plasticity
Changes in pore fluid chemistry
Fluid assisted diffusive mass transfer processes
Addition/production of melt
Example(s)
Horda Platform, Northern North Sea (Steel & Ryseth 1990); Brent
and Hutton oil-fields, North Sea (Yielding et al. 1991)
Loppa High, Norwegian Barents Sea (Wood et al. 1989);
East Shetland Basin northern North Sea (Roberts et al. 1995)
Broad Fourteens Basin, North Sea and others (Cooper & Williams
1989)
Masada, Dead Sea Graben (Marco & Agnon 1995)
Midcontinent fault and fold zones, USA (Barrs et al. 1995; Marshak
& Paulsen 1996); Yorkshire coalfield, UK. (Clausen et al. 1994)
‘Tectonic cyclothems’, various settings (Blair & Bilodeau 1988)
Gander–Avalon boundary, Newfoundland (Holdsworth 1994); Outer
Hebrides Fault Zone (Butler et al. 1995)
Ikertok shear belt and elsewhere (Grocott 1977); Outer Hebrides
Fault Zone (Sibson 1977; Butler et al. 1995)
Radiometric, Brevard Fault Zone USA (Fullagar 1992); Fission
track, Alpine Fault Zone, New Zealand (White & Green 1986);
Palaeomagnetic, Dalsfjord Fault, Norway (Torsvik et al. 1992)
Numerous examples (e.g. see Bartholomew et al. 1992; Stewart &
Hancock 1994)
Southern and eastern Africa and other areas (Sykes 1978; Daly et al.
1989)
‘Morphotectonics’ in Australia and elsewhere (Hills 1956); Quaternary faulting studies, various settings (Stewart & Hancock 1994)
synchronously with deformation and in the long-term (Table 2;
see reviews by White et al. 1986; Handy 1989; Rubie 1990).
Many are associated with the migration of hydrous fluids or
magmas (Davidson et al. 1994; Wintsch et al. 1995). It is likely,
therefore, that pre-existing faults and shear zones undergo
reactivation because they are weak (Hills 1956; White et al.
1986; Prucha 1992). This effect will be significant on a large
scale where faults or shear zones cut through the main
load-bearing regions of the lithosphere, i.e. the upper mantle
and middle crust in continental regions (Molnar 1988). Ultimately, a clearer understanding of the underlying controls of
reactivation should emerge from studies of fault and shear
zone rheology in ancient, exhumed regions where mid-crustal
and upper mantle rocks are exposed at the surface at present.
Conclusions
(1) Structural reactivation is a fundamental feature of deformation in the continental lithosphere. Old structures form
long-lived zones of weakness that tend to repeatedly accommodate successive crustal strains, often in preference to the
formation of new zones of displacement.
(2) An accurate picture of the role played by reactivation
during continental deformation will only emerge from the
rigorous application of criteria that demonstrate repetition of
displacement using absolute or relative time markers. The
indiscriminate use of geometric similarity as a reactivation
criterion needs to be avoided.
(3) Recurrent fault displacements that occur as part of the
seismic cycle will not normally be resolved by geological
RECOGNITION OF REACTIVATION
criteria in ancient settings. In neotectonic environments, additional reference to the current tectonic regime is necessary in
order to recognise reactivation.
(4) Reactivation occurs because fault and shear zone
processes often lead to weakening. These processes will be
especially important where faults cut through the main
load-bearing regions of the lithosphere.
The authors thank the participants of the London conference whose
comments have helped to shape some of the ideas expressed here.
Comments by R. Twiss, R. Musson, D. Peacock and especially
J. Walsh were also very helpful. R. Butler made sure the paper was
politically correct. Amerada Hess Ltd are thanked for funding the
work of C.A.B. and R.E.H. K. Atkinson is thanked for all her
expert assistance with the diagrams.
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Received 1 May 1996; revised typescript accepted 16 August 1996.
Scientific editing by Rob Butler and Alex Maltman.